Longitudinal contrast in turbulence along a ∼ 19° S section in the Pacific and its consequences for biogeochemical fluxes

Longitudinal contrast in turbulence in the Pacific

Longitudinal contrast in turbulence along a ∼ 19° S section in the Pacific and its consequences for biogeochemical fluxesLongitudinal contrast in turbulence in the PacificPascale Bouruet-Aubertot et al.

Microstructure measurements were performed along the
OUTPACE longitudinal transect in the tropical Pacific
(Moutin and Bonnet, 2015). Small-scale dynamics and turbulence in the first
800 m surface layer were characterized based on hydrographic and current
measurements at fine vertical scale and turbulence measurements at
centimeter scale using a vertical microstructure profiler. The possible
impact of turbulence on biogeochemical budgets in the surface layer was also
addressed in this region of increasing oligotrophy to the east. The
dissipation rate of turbulent kinetic energy, ϵ, showed an
interesting contrast along the longitudinal transect with stronger
turbulence in the west, i.e., the Melanesian Archipelago, compared to the
east, within the South Pacific Subtropical Gyre, with a variation of
ϵ by a factor of 3 within [100–500 m]. The layer with enhanced
turbulence decreased in vertical extent travelling eastward. This spatial
pattern was correlated with the energy level of the internal wave field,
higher in the west compared to the east. The difference in wave energy mostly
resulted from enhanced wind power input into inertial motions in the west.
Moreover, three long-duration stations were sampled along the cruise
transect, each over three inertial periods. The analysis from the western
long-duration station gave evidence of an energetic baroclinic near-inertial
wave that was responsible for the enhanced ϵ, observed within a
50–250 m layer, with a value of 8×10-9 W kg−1, about 8
times larger than at the eastern long-duration stations. Averaged nitrate
turbulent diffusive fluxes in a 100 m layer below the top of the
nitracline were about twice larger west of 170∘ W due to the
higher vertical diffusion coefficient. In the photic layer, the
depth-averaged nitrate turbulent diffusive flux strongly decreased
eastward, with an averaged value of 11 µmolm-2d-1 west of
170∘ W compared with the 3 µmolm-2d-1 averaged
value east of 170∘ W. Contrastingly, phosphate turbulent diffusive
fluxes were significantly larger in the photic layer. This input may have an
important role in sustaining the development of N2-fixing organisms
that were shown to be the main primary contributors to the biological pump in
the area. The time–space intermittency of mixing events, intrinsic to
turbulence, was underlined, but its consequences for micro-organisms would
deserve a dedicated study.

The subtropical South Pacific is one of the main oceanic deserts
characterized by an increasing oligotrophy to the east and the centre of the
gyre. A 43-day long cruise, the OUTPACE experiment, was performed in this
region, along an ∼1919∘ S longitudinal transect, during the
2015 austral summer in order to characterize the biological pump and its
coupling with dynamical processes (Moutin et al., 2017). In
addition to the trophic gradient the OUTPACE transect is also characterized
by a longitudinal contrast in dynamics between the “energetic” Melanesian
Archipelago (MA) and the “quiet” South Pacific Subtropical Gyre (SPSG)
(Rousselet et al., 2018). Hence the OUTPACE experiment
provides a unique opportunity to focus on physical and biological
interactions (Rousselet et al., 2018) that may prove crucial
in understanding biological pump functioning
(Ascani et al., 2013; Guidi et al., 2012). The influence of the mesoscale
and submesoscale circulations on the spatial distribution and transport was
detailed by Rousselet et al. (2018). In particular, they showed the strong
impact of fronts on the spatial distribution of bacteria and phytoplankton. A
detailed study of an anomalous surface bloom event by
de Verneil et al. (2017) revealed instead the main impact of
mesoscale advection. At smaller scales three-dimensional turbulence may have
a strong impact on the biological pump through the input of nutrients into
the photic layer and more generally in enhancing, in the stratified ocean,
vertical transports through turbulent diffusion
(Ledwell et al., 2008).

The level of turbulence is almost unknown in the OUTPACE area. To our
knowledge, the only microstructure measurements were performed in the western
part of the subtropical South Pacific during the Malaspina expedition
(Fernández-Castro et al., 2014, 2015) as part of an
extensive microstructure survey in the tropical and subtropical oceans. For
the leg done in the OUTPACE region, the averaged ϵ below the mixed
layer down to ∼300 m depth was ∼10-8 W kg−1, well
above the typical background dissipation rate for open ocean. Indirect
estimates of ϵ based on ARGO floats data fall in the same range as
Fernández-Castro et al. (2014) as shown by Whalen et al. (2012). This
study based on the global-scale ARGO floats dataset also revealed that the
southern subtropical Pacific is one of the most undersampled areas. At the
larger scale of the South Pacific Ocean, the equatorial zone is well known as
a hotspot for turbulence where shear instability prevails as a result of the
strongly sheared current system
(Gregg et al., 1985; Richards et al., 2015; Smyth and Moum, 2013; Sun et al., 1998). At
subtropical latitudes, where the background shear is lower, internal waves
are expected to play a major role in the onset of turbulence in the
stratified interior. Global maps of energy flux show enhanced semi-diurnal
tide energy conversion in the western part of the subtropical South Pacific
(Alford and Zhao, 2007a). The annual mean energy flux into
inertial motions is enhanced at mid-latitudes in all ocean basins with also a
south-east-oriented track in the Pacific from the Equator to 40∘ S
and within ∼180∘ E–160∘ W longitude in the OUTPACE
region (Alford and Zhao, 2007a). The latter process is subject
to seasonal variations, especially in subtropical regions where the
generation of energetic baroclinic near-inertial waves is favoured during the
cyclone season (Liu et al., 2008).

The contribution of the biological pump in the OUTPACE region to the main C,
N, and P biogeochemical cycles was one of the main purposes of the OUTPACE
project (Moutin et al., 2017). Moutin et al. (2018)
built a first-order budget at daily scale of these main elements, while
Caffin et al. (2018) focused on the role of N2 fixation.
N2 fixation was evidenced as the dominant process involved in the N
cycle in regions where Trichodesmium dominate. The input of nitrate
through turbulent diffusion was found to make a negligible contribution in
the photic layer as a result of a very deep nitracline. This raised the
question of the available source of other nutrients in the photic layer that
could sustain the development of N2-fixing organisms, the main
primary contributors to the biological pump in the area
(Caffin et al., 2018).

The purpose of this paper is to characterize the spatial variability of
turbulence along the OUTPACE transect with microstructure measurements
performed at both 1-day short-duration stations and at long-duration stations
lasting three inertial periods. The idea is also to provide insights into the
main mechanisms responsible for the observed turbulence with a focus on
long-duration stations that allow a characterization of the internal wave
field. How this small-scale dynamics influences biogeochemical fluxes is
another issue that is eventually addressed.

The OUTPACE cruise took place in early 2015 from 18 February to 3 April
onboard French oceanographic research vessel l'Atalante(Moutin et al., 2017). A set of 15 short-duration stations
(SD) over 24 h as well as 3 long-duration stations (LD) over three inertial
periods (the inertial period being ∼36 h) were performed along an
almost zonal transect starting from west of New Caledonia and ending near
Tahiti (Fig. 1).

2.1 CTD and LADCP

Conductivity–temperature–depth (CTD) measurements were performed on a
rosette using a SeaBird SBE 9plus instrument. Data were averaged over 1 m
bins to filter out spurious salinity peaks using Sea-Bird electronics
software. Simultaneously, currents were measured from a 300 kHz RDI lowered
broadband acoustic Doppler current profiler (LADCP). LADCP data were
processed using the Visbeck inversion method (Visbeck, 2002) and
provided vertical profiles of horizontal currents at 8 m resolution. These
measurements were performed at all stations with a typical 3 h time interval
between each deployment. In addition the ship was equipped with two SADCPs,
RDI Ocean Surveyors with frequencies 150 and 38 kHz yielding processed
currents averaged over 2 min time intervals and with vertical bins of 8 and
24 m respectively. Shear was computed using finite differences with current
vertical profiles interpolated over a 1 m vertical grid with an estimated
noise level of 5×10-4 s−1.

2.2 Microstructure measurements with VMP1000

Microstructure measurements were collected using a vertical microstructure
profiler, “VMP1000” (Rockland Scientific). This tethered profiler was
equipped with microstructure sensors, two shear sensors, and one
temperature sensor, as well as with Sea-Bird temperature and conductivity
sensors and a high-frequency fluorometer. A total number of 123 profiles were
performed with repeated profiles at LD stations (∼ 30 profiles over
three inertial periods) and at least one profile at each SD station except at
SD13 (see Table 1 for further details). The dissipation
rate of turbulent kinetic energy (ϵ) was inferred from
centimeter-scale shear
measurements. The vertical wavenumber shear spectrum was computed within the
inertial range, typically within meter to centimeter scales. The experimental spectrum was next compared to the
empirical spectrum, the Nasmyth spectrum (Nasmyth, 1970), which
allowed validation of the estimate of ϵ (Ferron et al., 2014). Shear measurements were processed using
the routines developed by Rockland Scientific. Specific noise removal
procedures were applied with the spikes in the shear data first removed and
spectral coherence between the shear sensors and the accelerometers used to
remove vibrational contamination. The first 20 m below the surface were not
considered to avoid any contamination from the ship wake as well as the 20 m
at the end of the profile because of the decreasing vertical velocity there.
More generally ϵ values were excluded when the vertical velocity
gradient of the VMP was larger than 2.5×10-2 s−1. The
averaged ϵ from the two shear probes was taken provided that the
ratio between the two estimates was smaller than 2; otherwise, the ϵ
value with the smallest depth variation (compared to the neighbouring upper
and lower ϵ values) was considered. ϵ was computed over a
1 m depth interval, and then a 8 m moving average was applied to this
signal. The estimated noise level is 5×10-11 W kg−1
following Ferron et al. (2014).

2.3 Diffusivity estimates

The diapycnal diffusivity, Kz, is commonly inferred from the kinetic
energy dissipation rate using the Osborn (1980) relationship:

(1)Kz=ΓϵN-2,

where Γ is a mixing efficiency defined as the ratio between the
buoyancy flux and the dissipation rate,
Γ=-gρ0ρ′w′‾ϵ, with w′ and
ρ′ the vertical velocity and density fluctuations, and N the buoyancy
frequency, inferred from the sorted density profile in order to avoid
spurious negative values associated with overturns,
N=-gρ0dρsorteddz,
with an 8 m moving average then applied to this signal. Γ was
generally set to 0.2 until the recent findings of Shih et al. (2005) and
Bouffard and Boegman (2013). These authors found a decrease in Γ for
increasing turbulence intensity, I, defined as

(2)I=ϵ/(νN2),

where ν is the molecular viscosity, ν=1.2×10-6 m2 s−1. In terms of timescales, I is the
ratio of the square of the Kolmogorov timescale, namely the dissipation
timescale of eddies at the Kolmogorov scale (ν/ϵ), and the
buoyancy timescale (1∕N). Shih et al. (2005) showed in a numerical study
that the Osborn relationship overestimated Kz when I>100 and proposed a
new parameterization of Kz for this regime. A few years later
Bouffard and Boegman (2013) proposed a refined parameterization of Kz
including in situ microstructure measurements in lakes as well. They defined
different regimes with the following formulations for Kz:

Kz=10-7 m2 s−1 within the diffusive sub-regime, I<1.7
;

Kz=0.171/4νI3/2 within the buoyancy-controlled sub-regime, I within
[1.7;8.5];

Kz=0.2νI, i.e., the Osborn relationship within the intermediate regime, I within
[8.5,400]; and

Kz=4νI1/2 within the energetic regime, I>400.

Note that for the OUTPACE dataset where most I values are smaller than 100,
the Kz values inferred from the Bouffard and Boegman parameterization that
is applied here do not differ significantly from those inferred from the
Osborn relationship.

The internal tide generation is inferred from the depth
integrated generating force following the linear approximation
(Baines, 1982) that reads as

‖F‖=∫N2ωz|Q⋅∇h|h2dz,

where N is inferred from the World Ocean Atlas monthly climatology,
ω is the tidal frequency, Q is the barotropic tidal flux, and
h is the bottom depth. The barotropic tidal flux is inferred from the
1/30∘×1/30∘ global inverse tidal model TPXO
(Egbert and Erofeeva, 2002) for two main constituents, the diurnal K1 and
the semi-diurnal M2.

The generation of baroclinic near-inertial waves occurs through inertial
pumping at the base of the mixed layer (Gill, 1984; Price, 1984).
Insight into possible generation of baroclinic near-inertial waves is
estimated from the wind work on inertial oscillations following
Alford (2003):

(3)Ff=ρτf2rH,

where τf2 is the square of the wind stress at the inertial
frequency, r is the damping of near-inertial motions in the mixed layer as
a result of baroclinic near-inertial wave radiation expressed as a function
of the inertial frequency r=0.15f, and H is the mixed-layer depth. The
wind stress was inferred from numerical simulations over the time period of
the cruise (Skamarock et al., 2005) and the mixed-layer depth from the seasonal
climatology.

2.5 Biogeochemical turbulent diffusive fluxes

The vertical components of nitrate and phosphate turbulent diffusive fluxes
were computed at all stations, using the diapycnal diffusivity, Kz,
inferred from microstructure measurements:

(4)FNO3,PO4=-Kz∂zcNO3,PO4,

where cNO3 and cPO4 are the nitrate and phosphate
concentrations. Measurements were performed daily at 12 depths in the first
200 m using standard colorimetric procedures Caffin et al. (2018). The quantification limits were
50 µmol m−3. At LD stations where six profiles were obtained,
the quantification limit for the mean concentration dropped to
20 µmol m−3 (i.e., 50/6). Concentration values below
the threshold for quantification were set to NaN. Vertical concentration
profiles were interpolated over a 1 m vertical grid and a 10 m moving
average was next applied. The top of the nitracline was defined as the depth
where nitrate concentration is zero based on an extrapolation from the last
detectable concentration (>50µmol m−3) assuming a
constant vertical gradient above this depth Moutin et al. (2018). It
is only at long-duration stations that the top of the nitracline was defined
in isopycnal coordinates: this allows us to get rid of a varying depth of the
nitracline because of vertical displacements of isopycnals induced by
internal waves Caffin et al. (2018). The minimum turbulent diffusive
flux was estimated from the threshold concentration with the molecular Kz
value (Kz∼10-7 m2 s−1):
0.4µmolm-2d-1 at SD stations and
0.17µmolm-2d-1 at LD stations where the mean
concentration profile was taken to compute the flux. A reference profile was
also defined in order to compare the nitrate diffusive flux variations along
the OUTPACE within a 100 m layer below the top of the nitracline. To do so
vertical profiles of concentration were first shifted vertically so that the
reference depth matched with the top of the nitracline and the mean vertical
cNO3 profile within the nitracline was inferred from this set of
re-scaled concentration profiles. Kz profiles were also rescaled onto this
vertical grid and a mean Kz profile inferred. The relative contributions
of the variations in Kz and that of ∂zcNO3 with respect to the total variations of the flux
were also inferred.

Figure 3Log values of the dissipation rate of turbulent kinetic energy
(W kg−1) (a) and vertical diffusion coefficient
(m2 s−1) (b) with N2−S2 in background color scale. The
number of profiles at each station is displayed with black circles and
profiles have been time-averaged at each station for better visualization.
The mixed-layer depth is plotted (black dashed curve) as well as the limit
between the two regions (dashed white line).

3.1 Overview

An overview of the spatial pattern of turbulence is given with depth-averaged
values of ϵ and Kz below 100 m depth at each station
(Fig. 1). Depth-averaged dissipation rates,
〈ϵ〉, vary within an order of magnitude within
[10-9.5;10-8.5] W kg−1. The highest values are observed west
of 170∘ W, in the shallower part, while the lowest values are
observed east of 170∘ W, in the deeper part. The same contrast is
retrieved on 〈Kz〉 with values ranging within
[10-5.8;10-4.8] m2 s−1. The western part of our study
area which shows the highest turbulence level is also the region where the
most intense velocities are observed, as illustrated with altimetry-derived
currents produced by AVISO along the RV l'Atalante cruise path
(Fig. 2a). The vertical section of the total velocity
modulus inferred from the SADCP data shows that this contrast is also
observed at depth with slightly larger velocities in the western part of our
study area (Fig. 2b). There the bathymetry ranges
typically from 4000 m up to a few hundred meters locally with significant
topographic slopes, which is consistent with the higher velocity signal; by
comparison, in the east the bathymetry is almost flat, with ∼5000 m
depth. More insights into turbulence are given with vertical sections of
ϵ and Kz in Fig. 3a and b. The range
of ϵ values covers 3 orders of magnitude,
∼[10-10,10-7] W kg−1 below the mixed layer (magenta curve
in Fig. 3a) down to 500 m depth. ϵ
presents a typical patchy pattern with spots of intense turbulence with
values up to ∼10-8 W kg−1 down to 500 m. These events are
localized over a 10 m vertical scale except at LD-A around 165∘ E
longitude where a 200 m layer of enhanced ϵ is observed. Most of
these events are observed in the west, west of 170∘ W, and the few
vertically localized large ϵ observed east of 170∘ W are no
deeper than 200 m depth (Fig. 3a). Statistics
of ϵ within a 100–500 m depth interval are given for each region
in Table 2. The contrast between the mean 100–500 m depth-averaged
ϵ in each of the two regions is a factor of 3 (Table 2). The
contrast is larger for the standard deviation, a factor of 10, which points
to the larger intermittency of turbulent events west of 170∘ W. The
Kz pattern presents a similar contrast between the western and eastern
parts with mean 100–500 m depth-averaged Kz between the two regions
differing by a factor of 2 (Table 2). More importantly, the standard
deviation of Kz is by far larger west of 170∘ W, by a factor of
15 (Table 2).

Figure 4Buoyancy frequency, N(a), shear, S(b), and
N2−S2(c) as a function of longitude and depth. The locations of
the CTD and LADCP profiles are displayed with black dashed lines. S is
inferred from LADCP data and N was vertically averaged over 8 m length
vertical intervals for consistency with the 8 m bin of the LADCP data.

Figure 5(a) Depth-integrated generation force for the K1 tidal constituent (log10(m2 s−2));
(b) same as in (a) but for the M2 tidal constituent. The
stations are shown with a red circle and the LD stations are indicated.

3.2 Shear instability

To gain further insights into the origin of this contrast in turbulence, the
possible occurrence of shear instability is addressed with buoyancy
frequency, N, shear, S, and N2−S2 sections displayed in
Fig. 4. The stratification is strong in the first
100 m, with a pycnocline that is generally well marked except westward of
165∘ E and around 172∘ W longitudes and a less stratified
surface layer above 50 m east of 170∘ W. The shear is significantly
higher west of 170∘ W, with high values over the 500 m depth
layer of the measurements in the west except for a 165–170∘ E
region (Fig. 4a and b). The likeliness of shear
instability was estimated from N2−S2, displayed in
Fig. 4c. The upper 100 m surface layer is the most
stable, with the highest N2−S2 values as a result of both strong
stratification and low shear. Shear instability is more likely to occur west
of 170∘ W below the 100 m surface layer of strong stratification
and where the shear is large: the percentage of data points that verifies the
criterion for shear instability is 10 times larger west of 170∘ W
than east of 170∘ W (Table 2). The fact that most of the subcritical
Ri, N2-S2<0, are observed west of 170∘ W is consistent with
the enhanced dissipation observed there (Fig. 3
and 4c).

3.3 Tidal and atmospheric forcings for the internal wave field

The possible impact of internal waves was estimated indirectly through the
two main energy sources of these waves, namely tidal forcing and wind power
input (Figs. 5 and 6).

The depth-integrated tidal generation force is displayed in
Fig. 5 for the K1 and M2 constituents. There is
a strong similarity between the two constituents with a generation that is
favoured in the western part which is shallower and with stronger topographic
gradients than the eastern part of our study area. The westernmost region is
characterized by numerous spots of generation with a depth-integrated
generation force of 103 m2 s−2, while eastward of
170∘ E longitude there is only one main generation site around
180∘ E longitude (Fig. 5a). This spatial
distribution of the internal tide forcing might suggest a similar contrast in
the internal-tide-induced dissipation since the high modes responsible for
turbulence are expected to dissipate within a few tens of kilometers of the
generation site (St Laurent et al., 2002).

Maps of wind power input on inertial motions, also referred to as inertial
flux, were computed using the spectral method described by
Alford (2003). The wind stress data were inferred from WRF numerical
simulations (Klemp et al., 2007; Skamarock et al., 2005) and the seasonal climatology
was used for the mixed-layer depth. The power input into inertial motions
gives insights into the generation of baroclinic near-inertial waves (niw) at
the base of the mixed layer through inertial pumping
(Gill, 1984). The maps reveal a strong longitudinal contrast in
inertial flux until mid-March (Fig. 6a–e). The
strongest wind power input was observed in the western part of our study
area. This is consistent with the climatology of storms and cyclones in the
area that are typically formed in the south-western tropical Pacific
Diamond et al. (2013). At the beginning of the cruise the largest
power input was localized south-west of the cruise stations
(Fig. 6a). Later a major event was observed
(Fig. 6d) during the passage of a tropical cyclone
over the area while RV l'Atalante was sampling to the east.
Eventually by the end of the cruise the inertial flux was small over the
OUTPACE region, with one spot of weak inertial flux observed in the east
(Fig. 6f–g). These maps suggest that energetic niw
are likely to be generated in the western part of our study area prior to the
cruise and until mid-March (Fig. 6a–b). The first
event of large inertial flux prior to the cruise may be particularly
insightful since it is likely to lead to the generation of equatorward
propagation niw within the OUTPACE sampling area, a scenario which is
consistent with large ϵ values there (Fig. 1).

Figure 6Maps of inertial energy flux (log10(W m−2)) every 7 days
during the cruise. Long-duration stations are shown with magenta
circles, and LD-A was held
during (b) and (c), LD-B during (e), and LD-C
during (f). The ship position is displayed with a magenta star.

Three long-duration stations were sampled each over three inertial periods,
LD-A in the western part of the transect and LD-B and LD-C in the eastern
part of the transect (see Table 1 and Fig. 1). Turbulence at LD-A
is by far the largest down to 500 m depth, with contrasted mean ϵ
and Kz between LD-A on the one hand and LD-B and LD-C on the other hand
(Fig. 7a and b) by factors of 7 and 5 for the
averaged, within [100 m, 500 m], ϵ and Kz respectively
(Table 3). Possible occurrence of shear instability is examined by comparing
mean profiles of shear square, S2, and N2 (Fig. 7c).
While the mean stratification is fairly close, at the three stations the
shear square is larger at LD-A compared to LD-B and LD-C within a factor of
10 within [50 m, 250 m] (Fig. 7c). Furthermore, within
[100 m, 200 m] S2 is larger than N2 at LD-A, pointing to possible
shear instability. This depth range coincides with local ϵ maxima,
thus reinforcing the shear instability hypothesis.

Figure 7Mean profiles at long-duration stations: ϵ in (a),
Kz in (b), and S2 and N2 in (c). Kz was
computed using N from the VMP measurements, while the profiles in
(c) were inferred from the rosette-mounted CTD and LADCP
instruments. The 95 % confidence interval for ϵ and Kz is
displayed with a shading surface in (a) and (b).

Figure 8Zonal velocity (m s−1) as a function of time and depth at the
long-duration stations; each row from 1 to 3 corresponds to LD-A, LD-B, and
LD-C respectively. In the first column ship ADCP data from the 150 kHz
instrument down to 300 m are displayed and in the second column those from
the 38 kHz instrument down to 800 m. Note that in (b) the 150 kHz
SADCP functioned only a few hours after the beginning of the station
sampling.

Figure 9Time-mean profiles of the kinetic energy of the subinertial flow,
the inertial and semi-diurnal frequencies, and ϵ at
LD-A (a), LD-B (b), and LD-C (c). The kinetic
energy was derived from the 38 kHz SADCP data but also from the 150 kHz
SADCP data for the inertial and semi-diurnal kinetic energies (thin blue and
red curves).

We next focused on the characterization of the internal wave field that may
reinforce the vertical shear and contribute to the onset of turbulence.
Current magnitude at LD-A is the largest (Fig. 8a and b), of
the order of 0.4 m s−1. Detailed insights from the 150 kHz SADCP
reveal a wavy pattern with two frequencies clearly identified
(Fig. 8a): strong upward-propagating bands close to the
inertial period, of about 1.5 days, and the semi-diurnal period, which
manifests itself through semi-diurnal heaving of upward-propagation niw
bands. The former is observed over the first 200 m only, while the latter is
observed down to ∼800 m depth (Fig. 8a and b). The
weaker currents at LD-B and LD-C are comparable with a maximum amplitude of
0.2 m s−1 (Fig. 8c–f). Periodic motions are also
evidenced with inertial oscillations in the first few meters
(Fig. 8c) and a combination of near-inertial and tidal
periods at depth. Noticeably, an upward phase propagation of niw can be
inferred at LD-B from the 38 kHz SADCP data (Fig. 8d). At
LD-C the semi-diurnal tidal signal dominates (Fig. 8f). The
dominance of niw at LD-A is consistent with the highest wind power input
inertial motions at LD-A (Fig. 6a) compared to LD-B
and LD-C (Fig. 6e). Instead the contrast observed
between the semi-diurnal depth-integrated generating force at LD-A compared
to that at LD-B and LD-C (Fig. 5b) is not evidenced on
the semi-diurnal currents (Fig. 8). This is well evidenced
below ∼300 m depth where the semi-diurnal tidal signal dominates at
all stations (Fig. 8b, d, and f). This difference might
result from localized generation areas of small scales that are not predicted
by the estimate performed with low-resolution fields (tidal model and
bathymetry) or from low modes with long-range propagation.

Table 3Statistics at the long-duration stations: percentage of N2-S2<0
(i.e., Ri<1), median values of ϵ, Kz, mean value of kinetic
energy for the sub-inertial frequencies, inertial frequency and semi-diurnal
M2 tidal constituent, and the same for shear variance. The average is
performed over a 100–500 m surface layer (first three lines) as well as
over a 50–250 m layer to highlight the impact of the niw at LD-A.

Figure 9 summarizes the main characteristics of the three
long-duration stations, with vertical profiles of time-averaged ϵ
and kinetic energy for the sub-inertial flow, the inertial frequency band,
and the semi-diurnal tidal constituent, M2. Kinetic energies were inferred
from the frequency spectra computed over three inertial periods.
Depth-averaged values of ϵ, Kz as well as kinetic energy and
shear within the different frequency bands are also given in Table 3. The
average was computed over two depth intervals: the first one is 100–500 m,
consistent with that chosen in Table 2, while the second one is 50–250 m,
corresponding to the depth interval of maximum niw energy at LD-A. The
enhanced ϵ at LD-A is coincident with an energetic niw
(Fig. 9a). The contrast with LD-B and LD-C is striking
within the depth interval 50–250 m with an averaged ϵ within
50–250 m about 8 times larger than at LD-B and a near-inertial kinetic
energy that differs by almost an order of magnitude (Table 3). The
significant decrease in ϵ, coincident with a sharp shutdown of the
near-inertial baroclinic signal around 250 m, shows the main effect of niw
on dissipation. This transition is associated with a strong variation in the
subinertial flow that suggests a wave mean flow interaction Soares et al. (2015). LD-B and LD-C present strong
similarities in ϵ and niw and M2 kinetic energies with 100–500 m
depth-averaged values that are close (Table 3). The local maxima in
near-inertial kinetic energy may evidence niw beams around 150, 450, and
650 m at LD-B and 200, 350, 400, and 650 m at LD-C
(Fig. 9b and c). The semi-diurnal kinetic energy presents
an interesting contrast between LD-B and LD-C: while it is larger at LD-B in
the first 500 m and smaller below, the opposite is observed at LD-C with
maximum semi-diurnal energy below 500 m depth, also suggesting a beam
structure. The subinertial flow is weakest at LD-C
(Fig. 9c) by a factor of 2 compared to LD-B and by a factor
of 8 compared to LD-A for the 100–500 m depth-averaged value of the kinetic
energy. Moreover, the depth-averaged low-frequency shear is below the noise
level at LD-B and LD-C (Table 3), both features suggesting a weak influence
on internal wave propagation. The subinertial flow is by far the largest at
LD-A down to ∼250 m (Fig. 9a). The contrast in
turbulence between the three stations is mostly confined in the upper few
hundred meters as a result of an energetic niw and its interaction with the
strongly sheared subinertial flow (Table 3). Deeper, variations in ϵ
and kinetic energies are much weaker and of the same order of magnitude at
the three stations (Fig. 9).

Figure 10Longitude–depth sections of chlorophyll concentration (a),
Kz(b), nitrate concentration (c), phosphate
concentration (d), nitrate turbulent diffusive flux (e),
and phosphate turbulent diffusive flux (f). The top of the
nitracline is shown with a magenta line, the depth of maximum chlorophyll concentration with a green
line, and the euphotic zone
depth with a dashed black line. The locations of the LD stations are shown
with a vertical dashed line. The scales of the color bar of the nitrate and phosphate concentration and turbulent diffusive
flux are set to allow a direct comparison of concentrations and fluxes
against nutrient requirements for phytoplankton (Redfield proportion:
N:P=16:1).

Table 5Relative contributions to the flux variations of Kz and cz
depth-averaged over a 100 m vertical layer defined from the top of the
nitracline, west of 170∘ W and east of 170∘ W and at
long-duration stations. The mean value and standard deviation are also
given.

5.1 Overview of the OUTPACE section

The distribution of chlorophyll concentrations along the OUTPACE transect
is typical of a transition from an oligotrophic area in the MA toward an
ultra-oligotrophic area in the SPSG (Moutin et al., 2017)
with a deepening of the deep chlorophyll maximum, DCM, from ∼60 to
∼160 m and an increase in the euphotic zone depth
(Fig. 10a). There is one noticeable exception
to this trend with a near-surface chlorophyll concentration maximum at
∼35 m depth, at LD-B. de Verneil et al. (2017), who focused on
the characterization of this anomalous phytoplankton bloom event, explained
its occurrence by the main impact of mesoscale advection and an island
effect. The deepening of the DCM results from that of the nitracline and from
the NO3 depletion in the first 200 m, an evolution consistent with
the increasing oligotrophy to the east
(Fig. 10a and c). Moreover, east of
172∘ W the nitracline is most often deeper than the base of the
euphotic layer, which reinforces the oligotrophy. The turbulent diffusive
nitrate flux displays the same longitudinal trend
(Fig. 10c) with a depth-averaged value in the
photic layer that differs by almost a factor of 4 west of 170∘ W and
east of 170∘ W (Table 4). The standard deviation is also by far
larger in the west, by a factor of 15. Whether these flux variations are
driven by Kz variations or cz variations was examined within a 100 m
depth interval below the nitracline all along the section. These large
variations of the turbulent diffusive nitrate flux at small scales
(Fig. 10c) mostly result from Kz
variations with a weaker contribution of variations in the vertical gradient
of nitrate concentration with a typical ratio of 3 between the two (Table 5).
The latter contribution tends to weaken slightly the eastward decrease in the
flux due to decreasing Kz. The variation of the nitrate diffusive flux was
also examined at LD stations only in order to focus on the impact of the
temporal intermittency of turbulent events at each LD station where a large
number of VMP profiles were collected. The flux variation resulting from
Kz variations, −ΔKzcz, and those resulting from cz
variations, −KzΔcz, were compared to the averaged flux over the
three LD stations. The −ΔKzcz always dominates and is by far the
largest at LD-A, with a relative value of 105 % compared with the
−50 % and −55 % relative values at LD-B and LD-C (Table 5). This
shows the major impact of the turbulence intermittency induced by the niw
event at LD-A.

Figure 11Time–depth sections of Kz(a, d, g), phosphate
turbulent diffusive flux (b, e, h), and nitrate turbulent diffusive
flux (c, f, i). The scales of the color bar
of the nitrate and phosphate diffusive flux are set to match with the typical
Redfield ratio in the area (N:P=16:1). The mean euphotic zone depth is
displayed with a black dashed line. The top of the nitracline is defined by
the isopycnal ρNO3 and falls at depths of ∼83.5 m at
LD-A, ∼111.2 m at LD-B, and ∼134.6 m at LD-C.

Phosphate turbulent diffusive flux is on average smaller than the nitrate
turbulent diffusive flux, but its relative impact may be rescaled by a
factor of 16 : 1 corresponding to the Redfield N : P ratio. Hence, for
visual comparison between the two fluxes, the scale for the phosphate
turbulent diffusive fluxes differs from that of the nitrate turbulent
diffusive flux within the Redfield ratio in
Fig. 10 and subsequently. The phosphate
turbulent diffusive flux displays a similar longitudinal gradient in the
first 200 m but presents an opposite trend in the photic layer, with a
depth-averaged value higher by a factor of 1.5 east of 170∘ W
(Table 5). The concentration of phosphate is not as strongly limiting as that
of nitrate in the photic layer with significant turbulent diffusive flux, in
a few spots, especially around 170 and 190 longitudes
(Fig. 10d). Consistently the same trend is
obtained for the phosphate concentration in the photic layer, with an
eastward increase in the photic layer as opposed to the nitrate concentration
that tends to zero at the eastern stations
(Fig. 10c and d). The absolute value of the
photic layer depth-averaged value of the flux, east of 170∘ W, is
4.01µmolm-2d-1. Considering a N : P Redfield ratio
of 16, the phosphate turbulent diffusive flux is significant compared to the
3.15µmolm-2d-1 value for the nitrate turbulent
diffusive flux (Fig. 10e and f, Table 4).
This striking difference in phosphate and nitrate turbulent diffusive fluxes
within the photic layer may play an important role in the development of
micro-organisms as discussed later.

Figure 12Histograms of Kz(a, c, e) and nitrate turbulent
diffusive flux (b, d, f) at long-duration stations LD-A, LD-B, and
LD-C around the top of the nitracline. The top of the nitracline is defined
by the isopycnal, ρNO3, with density values taken from
Caffin et al. (2018), Table 4. A density interval of the upper bound equal
to ρNO3+3×10-2 kg m−3, of a typical vertical extension equal to 3–4 m, was
chosen for the statistics. The mean value is shown in each subpanel with a dashed
line.

5.2 Focus on LD stations

Turbulent diffusive fluxes of nitrate and phosphate were further analysed at
long-duration stations (Fig. 11). The
time–depth evolution of Kz underlines the very large values encountered
at the most turbulent station, LD-A, in contrast with values at LD-B and LD-C
that show the occurrence of a few spots of intense mixing in a more quiescent
background with Kz∼3×10-6 m2 s−1
(Fig. 11a, d, and g). The largest nitrate
turbulent diffusive fluxes occur at LD-A
(Fig. 11b), while the smallest values are
observed at LD-B and LD-C (Fig. 11f
and i). The averaged values in the photic layer show that it is only at LD-A
that there is a small input of nitrate through diffusion with an average flux
of 8.41µmolm-2d-1 (Table 4).

In contrast, phosphate turbulent diffusive fluxes are significant well above
the euphotic zone depth at all LD stations
(Fig. 11b, e, and h), which may have an
impact on primary production, whereas the nitrate input by turbulent
diffusion is negligible (explanation below). Depth-averaged values in the
photic layer are even comparatively larger than that of the nitrate turbulent
diffusive fluxes at LD-A if one applies the P/N=1/16 Redfield ratio
(Table 4). Various spots of large phosphate turbulent diffusive fluxes are
also evidenced in the first ∼ 20–80 m that can be correlated with
events of intense turbulence (Fig. 11b, e,
and d). At LD-C the only event of significant phosphate turbulent diffusive
flux results from a strong turbulent event
(Fig. 11h).

5.2.1 Nitrate input at the top of the nitracline

The input of nitrate through turbulent diffusion was first examined at the
top of the nitracline, within a ∼4 m thickness layer, with histograms
of the nitrate turbulent diffusive flux (Fig. 12). There is a
strong contrast between LD-A and the eastern stations in the shape of the
histograms: an atypical shape of the histogram is obtained at LD-A with a
large standard deviation, while the distribution of the nitrate turbulent
diffusive flux is fairly similar at LD-B and LD-C with one main peak
(Fig. 12d and f). The atypical shape obtained at LD-A results from
the large intermittency of Kz as illustrated in Fig. 12a. The
distribution of Kz values covers 2 orders of magnitude and presents a
bimodal distribution. The moderate Kz values are associated with the
“background state”, while the large values are associated with intense
turbulent events related to the near-inertial baroclinic wave (e.g.,
Fig. 8a). The mean value of the nitrate turbulent diffusive
flux within a ∼4 m layer starting from the top of the nitracline is
smaller at LD-B compared to LD-A and LD-C: 24.1 and
18.9µmolm-2d-1 at LD-A and LD-C compared with the
mean value of 6.0µmolm-2d-1 at LD-B. This contrast
may appear surprising, with a comparable mean value at LD-A and LD-C, but
this results from the occurrence of a few turbulent events at LD-C where the
nitracline is deepest. Nevertheless, the main point regarding the impact of
the turbulent input of nitrate at the top of the nitracline on new primary
production is whether or not the top of the nitracline falls within the
photic layer. This point is addressed in the following with histograms in the
photic layer.

5.2.2 New primary production sustained by phosphate turbulent diffusive fluxes at the western stations?

Figure 13 summarizes the contrast between long-duration
stations in the photic layer with histograms of Kz, phosphate, and
nitrate turbulent diffusive fluxes. The euphotic zone depth (EZD) was
immediately determined on board from the photosynthetically available
radiation (PAR) at depth compared to the sea surface PAR(0+), and used to
determine the upper water sampling depths corresponding to 75, 54, 36, 19,
10, 3, 1 (EZD), 0.3, and 0.1 % of PAR(0+)
Herbland and Voituriez (1977); Moutin and Prieur (2012). The euphotic zone
depth varies within [55–120 m] with a mean value of 70 m at LD-A, 55 m at
LD-B, and 120 m at LD-C. As in the previous figures the scales for the
nitrate and phosphate turbulent diffusive fluxes match with the Redfield
ratio for visual comparison of the relative impact of each of these fluxes on
micro-organisms. Significant nitrate turbulent diffusive flux is observed at
LD-A as opposed to LD-B and LD-C as a result of shallower nitracline that
falls within the photic layer at LD-A (Fig. 13c, f, and i). At
LD-B and LD-C where the nitracline is below the euphotic zone depth the
nitrate turbulent diffusive flux is zero (Fig. 13f and i). The
station average of the phosphate turbulent diffusive flux is of the same
order of magnitude at LD-A and LD-B and smaller by a factor of 10 at LD-C
(Fig. 13b, e, and h; Table 4). These significant values of the
phosphate turbulent diffusive flux observed at LD-A and LD-B suggest an
impact on micro-organisms (Fig. 13e). Indeed, the presence of
nitrogen fixers Dupouy et al. (2000) and high rates of N input by N2
fixation were already noticed in the western tropical South Pacific
(Moutin et al., 2008) and confirmed in the whole area
(Bonnet et al., 2008). In contrast the smallest values observed at LD-C
(Fig. 13h) are consistent with low nitrogen fixation rates
measured there (Bonnet et al., 2017) or in the whole South Pacific gyre
(Moutin et al., 2008), probably because of iron depletion
(Blain et al., 2007; Bonnet et al., 2017; Guieu et al., 2018; Moutin et al., 2008). The lack
of iron for N2 fixers may explain their lower presence in the gyre
and consequently the relatively higher phosphate concentrations measured
there in the upper layer. Phosphate is not depleted by the N2 fixers
even with relatively low turbulent diffusive fluxes of phosphate from below.

Figure 14Time average of vertical profiles of nitrate and phosphate turbulent
diffusive fluxes at long-duration stations LD-A (a),
LD-B (b), and LD-C (c). The scales of the nitrate and
phosphate diffusive flux are set to match with the typical Redfield ratio
(N:P=16:1).

Figure 14 summarizes the main features of the turbulent
diffusive fluxes with time-averaged vertical profiles. The double x axis
for the nitrate and phosphate turbulent diffusive fluxes is scaled within a
Redfield ratio so that the fluxes are superimposed if they follow the
Redfield proportion. At depth, NO3 and PO4 fluxes follow the
Redfield proportion. Closer to the surface, NO3 flux decreased: at
LD-A there is a small input of nitrate in the photic layer, while at the
eastern stations the nitrate flux vanishes above the base of the euphotic
zone and reached zero before PO4 fluxes. These significant phosphate
fluxes at shallower depths may potentially be fuelling nitrogen fixation,
significant PO4 sources through turbulent diffusion that are likely
to provide the required conditions for the growth of N2 fixers in the
Melanesian Archipelago. Because phosphate availability likely sustains
N2 fixation and therefore N input by N2 fixation in the WTSP,
new primary production (Dugdale and Goering, 1967) might be sustained by new P (or
new N including N2 fixation) in the photic zone.

5.2.3 Turbulent diffusion and the oligotrophic to ultra-oligotrophic conditions encountered during the OUTPACE cruise

The decrease in nitrate turbulent diffusive flux eastward was found to be
consistent with the increasing oligotrophy and the deepening of the
nitracline (Caffin et al., 2018; Moutin et al., 2018). As a result the
nitrate input into the photic layer through turbulent diffusion was found to
provide only a subordinate contribution to the N budget with a
1 %–8 % contribution to the new N (Caffin et al., 2018). In the
Melanesian Archipelago, the input of nitrate into the photic layer represents
a small contribution during the stratified period (Caffin et al., 2018) as
well as on an annual timescale, with a 46 µmolm-2d-1
input by turbulent diffusion compared with a
642 µmolm-2d-1 input of N by N2 fixation
Moutin et al. (2018).

In a 100 m layer starting from the top of the nitracline, the mean nitrate
turbulent diffusive flux varies by a factor of 3 between the western LD
station LD-A with a flux of ∼45µmolm-2d-1 and
the eastern LD-B and LD-C stations. The mean value obtained at LD-A is
smaller than the average value obtained in the oligotrophic eastern Atlantic,
∼140µmolm-2d-1, by Lewis et al. (1986),
suggesting an increased oligotrophy in the Pacific. It is typically 1 order
of magnitude smaller than the values inferred further south, ∼30∘ S, by Stevens and Sutton (2012), where both larger Kz and
nitrate vertical gradient are responsible for a larger nitrate turbulent
diffusive flux. The comparison with Lewis et al. (1986) in the Atlantic
Ocean also highlights the ultra-oligotrophy of the Pacific Ocean in the gyre
compared to their counterpart in the Atlantic Ocean with diffusive fluxes at
least 1 order of magnitude smaller as shown at LD-B and LD-C. The low and
deep nitrate turbulent diffusive fluxes may not explain the higher primary
production and N2 fixation rates observed in the upper 0–40 m
Moutin et al. (2018).

The input of phosphate in the photic layer was also addressed as a possible
source for sustaining the development of N2-fixing organisms.
Phosphate turbulent diffusive fluxes mean values were significant in the
photic layer with the exception of the easternmost station. In all cases, a
few events of large fluxes driven by localized intense turbulent events were
identified. The large variations in the turbulent diffusive fluxes resulting
from the occurrence of strong turbulent events were thus underlined with a
focus on long-duration stations (Caffin et al., 2018). This
raised the question of the estimate of the turbulent input of nitrate in the
photic layer when establishing C, N, and P budgets as well as the impact of
turbulence intermittency on micro-organisms (Liccardo et al., 2013).

Variations within a factor of 10 of the depth-integrated ϵ were
observed along the OUTPACE transect. The largest ϵ observed in the
west compare well with the few measurements performed by
Fernández-Castro et al. (2014, 2015) in the area during
the Malaspina expedition. The range of values is comparable with 80 % of
ϵ values within [6×10-10;10-8] W kg−1 in the
first 300 m below the mixed layer for the Malaspina expedition and 82 %
for the OUTPACE ϵ within [30 m, 300 m] west of
180∘ longitude. Shear instability was evidenced as
one main process responsible for turbulence, which is a well-known mechanism
in the strongly sheared Pacific equatorial currents
(Richards et al., 2015; Smyth and Moum, 2013). Richards et al. (2012) also
mentioned the modulation of the turbulence level over a 3-year period with
different ENSO states in the western equatorial Pacific, with a maximum shear
during La Niña events compared to El Niño events. How this turbulence
cycle is relevant to the OUTPACE region would be an interesting point to
address with possible higher turbulence level provided that the OUTPACE
cruise took place during an El Niño event. Shear instability was found to
be more likely to occur in the western part of our study area, with most
critical Ri encountered there. This basic analysis thus explained the
contrast in dissipation observed along the transect. Whether the onset of
shear instability may be driven by the low-frequency flow or the internal
wave field could not be inferred from the short-duration stations, but
insights into the impact of internal waves were given with estimates of the
two main forcings for internal waves and the analysis of LD stations. The
main forcings of internal waves were found to vary significantly along the
19∘ S longitudinal transect, thus pointing out the possible impact
of internal waves on the contrast in energy dissipation. The most striking
factor was related to the atmospheric forcing, with the occurrence of
cyclones in the west leading to an energetic baroclinic near-inertial wave
field. This internal wave component was characterized at the western
long-duration station, LD-A, as well as its impact on energy dissipation.
These scenarios are typically encountered in tropical regions where
baroclinic near-inertial waves are known to contribute to energy dissipation
in the upper ocean (Cuypers et al., 2013; Soares and Richards, 2013). The
process of dissipation is often constrained by the mean subinertial flow with
ray focusing or critical levels depending on the spatial structure of the
flow (Soares et al., 2015; Whitt and Thomas, 2013). These mechanisms were
not addressed here, but will be the focus of a future study using
observations at the LD-A site. The context of the OUTPACE cruise with
significant niw generated by a cyclone is very specific, of general interest
for the studied region, where these meteorological phenomena are frequent at
the end of the summer. It hides the more continuous influence of internal
tides as a turbulence driver. Our measurements show only a slightly larger
semi-diurnal kinetic energy in the west at LD-A compared to LD-B and LD-C,
but suggest a larger contrast in shear variance.

The impact of turbulence on biogeochemical fluxes and trophic gradients was
estimated based on nitrate and phosphate turbulent diffusive fluxes. The
nitrate input into the photic layer through turbulent diffusion was found to
provide only a subordinate contribution to the N budget as a result of the
eastward decrease in nitrate turbulent diffusive fluxes and the deepening of
the nitracline. New production is mainly supported by N2 fixation and
is located in the western tropical South Pacific. The higher phosphate
turbulent fluxes compared to nitrate fluxes provide an excess P relative to
Redfield stoichiometry and a potential ecological niche for N2 fixing
organisms in the west. In addition to the lower iron availability in the east
preventing N2 fixation from occurring, the high iron availability in
the west allows this process and the excess P provided to the photic zone to
sustain a higher new primary production, explaining the western–eastern
oligotrophy gradient.

This article is part of the special issue “Interactions between
planktonic organisms and biogeochemical cycles across trophic and N2
fixation gradients in the western tropical South Pacific Ocean: a
multidisciplinary approach (OUTPACE experiment)”. It is not associated with
a conference.

This is a contribution of the OUTPACE (Oligotrophy from
UlTra-Oligotrophy PAcific
Experiment) project (https://outpace.mio.univ-amu.fr, last access: 12
December 2018) funded by the French research national agency
(ANR-14-CE01-0007-01), the LEFE-CyBER programme (CNRS-INSU), the GOPS
programme (IRD), and the CNES (BC T23, ZBC 45000048836). We warmly
acknowledge the assistance of the crew of the French research vessel
l'Atalante during the deployment of the VMP. We thank
Olivier Desprez for his technical support and help during the cruise. The
microstructure profiler was funded through the ANR OUTPACE
project.

Moutin, T. and Prieur, L.: Influence of anticyclonic eddies on the
Biogeochemistry from the Oligotrophic to the Ultraoligotrophic Mediterranean
(BOUM cruise), Biogeosciences, 9, 3827–3855,
https://doi.org/10.5194/bg-9-3827-2012, 2012. a

The OUTPACE cruise took place between New Caledonia and French Polynesia. The main purpose was to understand how micro-organisms can survive in a very poor environment. One main source of nutrients is at depth, below the euphotic layer where micro-organisms live. The purpose of the turbulence measurements was to determine to which extent turbulence may uplift nutrients into the euphotic layer. The origin of the turbulence that was found contrasted along the transect was also determined.

The OUTPACE cruise took place between New Caledonia and French Polynesia. The main purpose was...